Marine Geology 286 (2011) 95–105
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Marine Geology j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / m a r g e o
Slope channel formation, evolution and backﬁlling in a wide shelf, passive continental margin (Northeastern Sardinia slope, Central Tyrrhenian Sea) Giacomo Dalla Valle ⁎, Fabiano Gamberi Istituto di Scienze Marine (ISMAR-Geologia Marina), CNR, via Gobetti 101, 40129, Bologna, Italy
a r t i c l e
i n f o
Article history: Received 10 December 2010 Received in revised form 24 May 2011 Accepted 9 June 2011 Available online 16 June 2011 Communicated by D.J.W. Piper Keywords: internal levee inner thalweg terrace knickpoint sea level variations channel backﬁlling
a b s t r a c t The sedimentary processes of the Caprera slope channel system (CSCS), located in the northern slope sector of the eastern Sardinian margin, have been reconstructed through the integrated interpretation of multibeam bathymetric data and seismic proﬁles. The CSCS is composed of three main slope channels (C0, C1, C2), and of a series of smaller V-shaped channel-forms, that show different morphologies and sedimentary inﬁll. The southernmost C0 slope channel, is abandoned and its upper segment is completely buried and healed, whereas its lower segment is still evident in the bathymetry. On the contrary, the upper segments of the northernmost slope channels, C1 and C2, are characterized by an inner thalweg, erosional terraces, and internal levees. The channel bends feature complex erosive geomorphic elements, mainly controlled by the interaction of sediment gravity ﬂows with the channel shape. The lower reaches of C1 and C2 slope channels have a ﬂat ﬂoor, which is the result of the deposition of channel-wide sedimentary bodies. The present setting of the CSCS is indicative of a waning phase and a decreased volume of sediment gravity ﬂows following a main waxing phase of channel excavation. This scenario can be interpreted as resulting from the decreased energy ﬂow following the last glacial maximum. During low stand periods, river-fed sediment gravity ﬂows had high energy and excavated the slope channels. During the present-day high stand, the main ﬂows, resulting from the interaction between alongshore and cross-shelf currents had a downbasin depositional behavior and resulted in the complete healing, or partial backﬁlling of the slope channels. © 2011 Elsevier B.V. All rights reserved.
1. Introduction Within submarine continental slopes, deep-water channels, i.e. those channels that connect the shelf-edge to the base of the continental slope and beyond, are prominent features that act as transport pathways and deliver sediment to the deep-water realm (Mayall and Stewart, 2000; Peakall et al., 2000; Posamentier and Kolla, 2003; Gee and Gawthorpe, 2007; Kane et al., 2007; Kolla, 2007; Schwarz and Arnott, 2007; Wynn et al., 2007; Gamberi and Rovere, 2011; Jobe et al., 2011; McHargue et al., 2011). In recent years, detailed studies of outcrops, modern systems, and experimental and laboratory works, also in comparison with ﬂuvial systems, have focused on the response of the slope channel architectural styles to environmental modiﬁcations driven by allogenic and autogenic controlling factors (Reading and Richards, 1994; Pirmez et al., 2000; Kneller, 2003; Bouma, 2004; Estrada et al., 2005; Gamberi and Marani, 2007; Kane et al., 2007; Peakall et al., 2007; Normark et al., 2009a,b; Dalla Valle and Gamberi, 2010; McHargue et al., 2011; Sylvester et al., 2011). Allogenic drivers, such as the hinterland geology and climate ﬂuctuations, impact on the shape, the morphology and the nature of the sediment source, and affect sediment production and transportation ⁎ Corresponding author. E-mail address: [email protected]
(G. Dalla Valle). 0025-3227/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.margeo.2011.06.005
rates and control the location and spacing of the ﬂuvial entry points in the coastal areas (Posamentier and Kolla, 2003; Bouma, 2004; Ducassou et al., 2009; Gamberi and Dalla Valle, 2009; Covault and Graham, 2010; Covault et al., 2010; Sømme et al., 2011). Sea level variations, through the modiﬁcation of the width and morphology of the continental shelf, control the amount of sediment stored in the shallow water areas with respect to that discharged into the deep-water environment (Shanley and McCabe, 1998; Goodbred, 2003; Bouma, 2004; Weimer and Slatt, 2004). Sea level variations can also inﬂuence the pathways and the intensity of longshore currents (Covault et al., 2007; Boyd et al., 2008; Romans et al., 2009). The combination of sea level variations and longshore currents controls both the rate and site of sediment deposition,which can lead to the activation or the demise of sedimentary routes to the deep-water, with the progressive downslope inﬁll of canyons or slope channels (Walsh et al., 2007; Jobe et al., 2011). In addition, contrary to sequence stratigraphy models, in canyon-slope channel systems, sediment gravity ﬂow can also be active during sea level highstands as documented in various modern and ancient turbidite systems (Khripounoff et al., 2003; Covault et al., 2007; Boyd et al., 2008; Carvajal and Steel, 2009; Normark et al., 2009a,b). The maintenance of channels during highstand depends largely on the geodynamic setting and the position of the slope channel head with respect to the sediment entry points, and on the width of the shelf (Covault et al., 2007).
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sinuous and meandering channels. Finally, we argue that the different evolutionary histories of each slope channel of the CSCS can be connected with in the effects of sea level variations on their source areas and, in turn, on the processes that initiate sediment gravity ﬂows.
The analysis of intraslope channel morphologies and internal geometries can give information on the channel history and on its response to variations in allogenic controlling factors such as sea level variations (Deptuck et al., 2007; Jobe et al., 2011; McHargue et al., 2011). In particular, many works have dealt with the understanding of the occurrence or absence of meander point bars within sinuous slope channels, as a function of the sediment gravity ﬂow behavior, grain size and trigger mechanisms (Abreu et al., 2003; Deptuck et al., 2003; Dykstra and Kneller, 2009; Khan and Arnott, 2011). Detailed observations on the geomorphic element along submarine channel bends, carried out both on modent systems, outcrop and subsurface data, are still relatively rare and not unambiguous (Fildani et al., 2006; Arnott, 2007; Peakall et al., 2007; Straub et al., 2008). Furthermore, experimental works have furnished considerable discrepancy in their results regarding ﬂow behavior in channel bends and their resultant internal architecture (Kassem and Imran, 2004; Peakall et al., 2007). Here, we present an example of the high variability of internal architectural elements within three channels that make up the Caprera slope channel system (CSCS) in the northern sector of the continental slope of the Olbia Basin (eastern Sardinian margin). Through the analysis of high resolution multibeam bathymetric data and Air Gun/ Sparker seismics we can interpret the sedimentary processes that are connected with the development of the slope channels in response to allogenic controlling factors as such sea level and alongshore current variations. We also focus on the analysis of the geomorphic elements found along the channel bends, and highlight that ﬂow behavior can occur substantially different from those general assumed to develop in
2. Regional setting The eastern Sardinian margin is the passive margin of the Tyrrhenian Sea back-arc basin (Gamberi and Marani, 2004) (Fig. 1). In the Sardinian sector, the rifting processes that led to the formation of the Tyrrhenian Sea, started in the Late Tortonian and were completed in the Early Pliocene (Kastens and Mascle, 1990; Sartori, 1990). Along the eastern Sardinian margin, the rift morphology is still evident in a series of inherited structural highs that bound the intraslope basins of the upper sector of the margin (Fig. 1). The Olbia Basin (OB) is the northernmost intraslope basin of the eastern Sardinian margin; it has a 25 km wide basin plain, located at around 1500 m of depth, and is limited seaward by the Baronie and the Etruschi seamounts (Gamberi and Marani, 2004; Dalla Valle and Gamberi, 2010) (Fig. 1). The continental slope of the OB is around 25 km wide, with the toe of the slope located at a depth of around 1200 m. It is ﬂanked landward by a 20 km wide shelf, with the shelfbreak located at a depth of around 120 m (Ulzega, 1987) (Fig. 1). Above the Messinian erosional surface, corresponding with a prominent reﬂector, the Pliocene–Quaternary sedimentary cover of the Olbia continental slope is around 0.25 s (t.w.t) thick, corresponding approximately to 220 m of sediments.
LSC Shelf break
Baronie seamount 11°
Tyrrhenian back-arc basin
ESM 25 km
Fig. 1. Shaded relief map of the northeastern Sardinian margin in the Tyrrhenian Sea from multibeam bathymetric data (see location in the lower right inset). The study area corresponds with the box and is shown in more detail in Fig. 2. The Caprera slope channel system (CSCS) located in the study area feeds the Caprera deep-sea fan (CDF) in the Olbia Basin plain that is bounded seaward by the Baronie and the Etruschi Seamount. The other slope channel systems of the Olbia Basin as the Mortorio (MSC) and the Tavolara (TSC) are also indicated. The shelf-break (dashed line) is taken from Ulzega (1987). LSC: southward longshore currents from Corsica; BSC: wind-driven currents from the Bonifacio Strait (from Astraldi and Gasperini, 1992 and Artale et al., 1994).
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The Caprera slope channel system (CSCS) is part of the main turbidite system of the northern sector of the Olbia continental slope that also comprises the Mortorio and the Tavolara turbidite systems (Figs. 1, 2) (Gamberi and Marani, 2004; Gamberi and Dalla Valle, 2009; Dalla Valle and Gamberi, 2010). The CSCS consists of three slope channels, here named C0, C1 and C2 (Figs. 1, 2). In the middle slope, the C1 and C2 channels join into a single trunk, named the Caprera slope channel which at the base of slope, feeds the Caprera deep sea fan (Dalla Valle and Gamberi, 2010) (Figs. 1, 2). 3. Data and methods The present study is based on the interpretation of data collected during the TIR-99 cruise carried out by ISMAR (Institute for Marine Science) of Bologna in 1999. A bathymetric survey was performed along the eastern Sardinian margin downslope at a depth of 500 m. The bathymetric data were acquired with a Kongsberg-Simrad EM12-120S and processed at ISMAR with Hips & Sips software by Caris (http:// www.caris.com). Bathymetric and shaded relief maps were obtained by gridding the data at intervals of 25/50 m using the Global Mapper software (http://www.globalmapper.com). Subsurface observations are allowed by a closely spaced grid of Air Gun proﬁles, collected during the multibeam survey, and a low density grid of 30 kJ Sparker proﬁles collected by ISMAR in the 70s during the BS-77 cruise (Fig. 2). 4. Results 4.1. The C0 slope channel The C0 is the southernmost channel of the CSCS, and consists of an upper, completely buried segment, and a lower, partially buried segment (Figs. 2, 3).
4.1.2. Lower segment The lower segment of the C0 slope channel, from around 700 m of depth to the base of the slope, is only partially buried, and its course is clearly evident in the bathymetric data (Fig. 3). This segment of the C0 is moderately sinuous (sinuosity index: 1.16), around 1 km wide, and is around 12 km long. The channel ﬂanks have heights ranging from 150 m (northern ﬂank) to around 100 m (southern ﬂank), and both lose relief downchannel (Fig. 3). The ﬂanks are markedly asymmetric,
C1 slope channel C2 slope channel
4.1.1. Upper segment A 4 km wide, downslope narrowing bathymetric low is the seaﬂoor expression of the buried segment of the C0 channel (Figs. 2, 3). Within the main bathymetric low, two subtle, linear C0a and C0b depressions are separated by a 4 km long ridge (Fig. 3). The C0a and C0b depressions have a SW–NE trend and converge at around 700 m of depth, just upslope from the lower reach of the C0 channel (Fig. 3). The two depressions have broad U-shaped ﬂat ﬂoors and are 15 m to 35 m deeper than the surrounding seaﬂoor, and lose relief downslope (Fig. 3). A seismic proﬁle shows that the C0a and C0b depressions are the seaﬂoor expression of two channel-forms buried below around 80 ms of sediments (~70 m) (Fig. 4a). The channel-form below the C0a depression, is around 800 m wide and has a U-shaped basal incisional surface, and is ﬁlled by a package of high-amplitude, parallel reﬂectors (Fig. 4a). The channel-form below the C0b depression has a V-shaped cross-section, and is ﬁlled with more irregular reﬂections sometimes displaying a cut-and ﬁll stacking pattern (Fig. 4a). To the north of the two depressions, a series of smaller, V-shaped channelforms (less than 500 m of width), ﬁlled with chaotic seismic facies, spans a belt of around 3.5 km (Fig. 4a). Similarly to the channel-forms corresponding with the C0a and C0b depressions, they are buried beneath parallel reﬂectors (Fig. 4a). In the seaﬂoor sector above the small V-shaped channel-forms, a series of crater-like depression, interpreted as pockmarks, is also present (Fig. 3).
C0 slope channel
Olbia continental slope
Caprera slope channel Base of slope
Fig. 2. 3D shaded-relief map of the northern sector of the Olbia slope, showing the C0, C1 and C2 channels that compose the Caprera slope channel system. The C1 and C2 slope channels converge into the Caprera slope channel, that at the base of slope feeds the Caprera leveed channel. The available seismic lines (Air Gun, dashed lines; Sparker, bold lines) are shown in the inset on the upper left (contour interval is 50 m).
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C1 slope channel H1 high
Base of the slope C0 lower segment C0a depression
Fig.4b C0 upper segment
C0b depression Fig.4a
Fig. 3. Shaded-relief map of the C0 slope channel and of the surrounding Olbia continental slope, showing the buried upper reach of the C0, with the C0a and C0b depressions and the partially buried lower reach. The northward ﬂanking H1 high is indicated. Pockmarks are mainly scattered above the buried reach of the C0 slope channel and on the top of the H1 high. The bold dashed black lines correspond to the traces of the seismic proﬁles of Fig. 4a and b.
From the upslope limit of our bathymetric data at a depth of 700 m, (down to the junction with the C2 channel) the C1 slope channel is around 17 km long (Figs. 2, 5). The C1 slope channel consists of two distinct, linear sectors separated by a sector with two turns from 850 m to 900 m in depth (Figs. 2, 5). The limit between the upper and the lower channel segments roughly corresponds with the second turn (Fig. 5).
thalweg course, and correspond to an increase of its gradient up to 7° (Figs. 5, 7a). The seismic proﬁle crossing this portion of the channel shows that the superﬁcial reﬂectors in the terraces are truncated by the thalweg incision (Fig. 6a). Thus, it is suggested that in this sector, the terraces represent intrachannel erosional remnants formed in consequence of the erosional activity of the sedimentary ﬂows that has carved the inner thalweg. The inner thalweg disappears downslope from the second knickpoint, at a depth of 810 m, where the channel ﬂoor becomes featureless for a brief tract (Figs. 5, 7a). In correspondence with the ﬁrst bend of the channel a knickpoint (Knickpoint A in Figs. 5, 7, 8a) marks the reappearance of the inner thalweg with the ﬂanking erosional terraces (Fig. 5). Knickpoint A corresponds to a 55 m high step along the thalweg ﬂoor and is followed downslope by a megascour about 1 km long ending against the channel master ﬂank (Figs. 7a, 8a). The inner thalweg is around 500 m wide, and is ﬂanked on both sides by erosional terraces that elevate by up to 30 m with respect to the thalweg ﬂoor (Figs. 5, 8a). An inner thalweg can be observed along the channel tract between the ﬁrst and the secondbends of the C1 channel (Fig. 5). In correspondence with the distal part of the second bend, the relief of the inner thalweg becomes progressively more subdued.
4.2.1. Upper segment The upper segment of the C1 channel is straight, around 1 km wide, bounded by scalloped, steep (up to 15°), 200 m high master ﬂanks (Fig. 5). Seismic proﬁles show evidences of long-term sediment instability processes along the channel master ﬂanks, as conﬁrmed by truncated reﬂectors and buried removal surfaces, and by glide planes that slightly lower the left channel ﬂank (Fig. 6a,b). An inner thalweg is present within the channel ﬂoor, and is bounded by 30 m high, intra-channel terraces (Fig. 5). The inner thalweg is, in general, centered within the channel ﬂoor although tracts where it is adjacent to one of the channel ﬂanks are also present (Figs. 5, 6a). It is around 100 m wide, broadens downslope, and has an average gradient of around 1° (Figs. 5, 7a). Two small knickpoints, at depths of 730 and 810 m respectively, are present along the inner
4.2.2. Lower segment The lower segment of the C1 slope channel is characterized by a 1 km wide, U-shaped ﬂoor, conﬁned by channel master ﬂanks, around 125 m high that progressively lose relief to around 100 m (Fig. 5). The channel ﬂoor is generally featureless, with the stretch between 900 and 920 m characterized by a gradient of around 0.3°; whereas the distal stretch of the lower segment has a gradient of 0.8° (Fig. 7a). In its proximal part, on the outside of the second channel bend, a triangular-shaped area, around 2 km long and 100 m high, is developed adjacent to its northern ﬂank (Fig. 5). It consists of upward convex reﬂectors that onlap against the channel master ﬂank (Fig. 6c). Therefore, this feature can be interpreted as an internal levee. The C1 slope channel is inﬁlled with at least 0.2 s (~180 m) of channel-wide, high-amplitude, continuous horizontal reﬂectors (HARs) (Fig. 6c,d).
both in relief and steepness, due to the presence, to the north of the channel, of the H1 bathymetric high (Figs. 3, 4b). The lower segment of the C0 slope channel is ﬂat bottomed, and is only partially inﬁlled. The basal inﬁll consists of high-amplitude, continuous, horizontal, parallel reﬂectors (HARs), that can be interpreted as the result of the channel-wide deposition of possibly sand-prone sediment (Fig. 4b). This unit is cut by inclined, and more discontinuous reﬂectors, that can be interpreted as slump deposits due to the failure of the channel ﬂanks (Fig. 4b). At the base of the slope, the C0 channel continues as a very subtle U-shaped depression parallel to the continental slope (Figs. 2, 3). 4.2. The C1 slope channel
G. Dalla Valle, F. Gamberi / Marine Geology 286 (2011) 95–105
Lobe Pliocene-Quaternary sediments
C0 lower segment
C0 Channel infill
Fig. 4. a) NW–SE trending Sparker seismic proﬁle of the upper segment of the C0 slope channel showing the two buried C0a and C0b channels below subtle depressions at the seaﬂoor and further small V-shaped channel-forms to the north (see Fig. 3 for the location of the proﬁle). b) SE–NW trending Sparker seismic proﬁle of lower Olbia continental slope showing the sedimentary inﬁll of the lower segment of the C0 slope channel (see Fig. 3 for the location of the proﬁle).
The master ﬂanks of the slope channel do not show evidence of signiﬁcant instability processes (Fig. 6c,d). 4.3. The C2 slope channel The C2 slope channel is the northernmost branch of the CSCS and consists of two proximal tributaries (C2a and C2b) that join downslope into a single element (Fig. 5). The C2 slope channel is around 16 km long, with straight segments, separated by two turns at a depth of 900 and 950 m (Fig. 5). 4.3.1. C2a and C2b tributaries Only the very distal part of the C2a and C2b tributaries are imaged in the available bathymetric data. The C2a is characterized by a deep V-shaped incision, lacking any inﬁll, suggesting that, at present, the tributary is dominated mainly by erosional ﬂows (Fig. 6a). In contrast, the C2b has a ﬂat ﬂoor with a basal thin lense of chaotic inﬁll. Its upper inﬁll consists of parallel low-amplitude reﬂectors that can be interpreted as hemipelagic sediments. This setting of the C2b inﬁll points to a depositional phase within this tributary channel (Fig. 6a). On both C2a and C2b ﬂanks abundant evidence of sediment failures is visible and thus we interpret the chaotic inﬁll of the C2b has resulted from slump deposits fed within the channel from ﬂank failures.
4.3.2. Upper segment The upper segment of the C2 slope channel has an average gradient of 1°, is very narrow and has a V-shaped inner thalweg that is less than 100 m wide (Fig. 5). A knickpoint is present at a depth of around 820 m corresponding to an increase in the channel axis gradient to around 8° (Fig. 5, 7b). To the north, the thalweg is ﬂanked by a linear terrace of around 4 km long with a relief of up to 100 m above the thalweg ﬂoor. The linear terrace is bordered by the 60 m high scarp of the northern master ﬂank of the channel (Fig. 5). The terrace has a markedly different seismic facies with respect to the bounding continental slope; in addition, its upper portion consists of convex-up, bedded reﬂectors that are plastered against the more irregular slope reﬂectors (Fig. 6b). Given its morphology and seismic character, this terrace is interpreted as an internal levee. 4.3.3. Lower segment In correspondence with the turns, the channel shows an abrupt knickpoint, followed downchannel by a tract with a gradient near to 0° (Fig. 7b). This low gradient area continues for around 1.5 km, down to a depth of 960 m and corresponds with an elongated spoon-shaped depression along the channel course (Figs. 5, 7b, 8b). This depression is bounded downchannel by a rampart elevated around 7 m above the channel ﬂoor (Figs. 5, 7b, 8b).
G. Dalla Valle, F. Gamberi / Marine Geology 286 (2011) 95–105 Fig. 8b
Fig. 6d Featureless channel floor
C2a Limit between the upper and the lower reaches
Fig. 6c Terraces Terraces Knickpoint A Scalloped flanks Fig. 6b Knickpoints
Fig. 5. Shaded-relief map of the C1 and C2 Caprera slope channels with their main sedimentary architectural elements. Terraces are developed along most of the upper reach of the C1 where also an inner thalweg is present. The terrace located in the outer bend of the ﬁrst turn of the C1 is ﬂat-topped, whereas the terrace located in the outer-bend of its second turn is articulated in a series of steps. Internal levees are presents in both slope channels. The limit between the upper and the lower reaches of the two slope channels is also indicated. The black dotted lines correspond with the traces of the seismic proﬁles of Fig. 6a, b, c, d.
Beyond the turns, the C2 channel enlarges progressively, reaching a width of about 700 m, and has a U-shaped ﬂat ﬂoor (Fig. 5), bounded by low relief ﬂanks (75–100 m high). At the junction with the C1 channel, the ﬂoor of the C2 slope channel is around 50 m deeper than the ﬂoor of the C1 channel (Fig. 5). A crossing seismic proﬁle shows that the lower segment of the C2 channel is inﬁlled with highamplitude, parallel, continuous reﬂectors (HARs) (Fig. 6d).
5. Discussion 5.1. Sediment gravity ﬂows and intrachannel elements In passive margins, the degree of activity and the erosional, bypassing or depositional character of the slope channels is in general, controlled by relative sea level variations. Sea level determines whether channel heads are connected with river mouths or to other feeding systems such as littoral drift cells or alongshore currents (Normark et al., 1998; Covault et al., 2007; Boyd et al., 2008). Furthermore, the behavior of slope channels can be inﬂuenced by changes in gradient due to variations in the geometry and the morphology of the receiving basin as a consequence of tectonic activity or sediment spill to a basin located at a deeper level (Pirmez et al., 2000; Sinclair and Tomasso, 2002; Gamberi and Marani, 2007; Dalla Valle and Gamberi, 2010). In the CSCS, only the C2a trunk, one of the tributaries to the C2 slope channel, is at present completely dominated by erosional processes as shown by its V-shaped proﬁle and by the lack of sedimentary inﬁll (Figs. 5 and 6a). On the contrary, the C1 and C2 and C2b slope channels show features that point to a present-day combination of erosional and depositional processes.
In the upper segments of the C1 and C2 slope channels, erosional processes are conﬁrmed by the incision of an inner thalweg, resulting in ﬂanking, marginal erosional terraces (Fig. 5). However, also deposition occurs within C1 and C2 upper segments resulting in the formation of intrachannel architectural elements, as internal levees conﬁned within the master ﬂanks (Fig. 5). In general, inner thalwegs and internal levees formed within the ﬂoor of larger valleys are recognized as an indication of intrachannel deposition due to small sediment gravity ﬂows that follow larger ﬂows responsible for the excavation of the main channel valley (Kneller, 1995; Torres et al., 1997; Weber et al., 1997; Piper et al., 1999; Mosher et al., 2004; Kane et al., 2007; Mosher and Piper, 2007; Kane and Hodgson, 2011). The lower reaches of C2 and C1 slope channels are, at present, the loci of deposition for sedimentary ﬂows within the slope channels as shown by their ﬂat bottom with the absence of any erosional feature (Fig. 5) and by their inﬁll consisting of continuous horizontal HARs (Fig. 6c,d). In the study area, as demonstrated by Dalla Valle and Gamberi (2010), during the present high-stand, sedimentary ﬂows are unable to reach the base of the slope and the Caprera deep sea fan (Fig. 1). We therefore assume that the C1 and C2 slope channels are experiencing a phase of transition from a waxing phase of the sedimentary ﬂows, which are responsible for the main channel valley excavation, to a waning phase of the sedimentary ﬂows (see Deptuck et al., 2007; McHargue et al., 2011). Erosional processes and the bypass of sediments are dominant in the upper reaches of the C1 and C2 channels, as the inner thalweg and bounding erosional terraces conﬁrm. On the contrary, in the lower reaches of the C1 and C2 slope channels, processes of channel backﬁlling, with the aggradation of possibly sand-rich deposits, conﬁned within the master ﬂanks of C1 and C2 channel, occur.
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d Fig. 6. a) NNW–SSE Sparker seismic proﬁle crossing the upper segment of the C1 slope channel and the C2a and C2b, tributaries to the C2 channel. b) N–S trending Air Gun seismic proﬁle of the upper segments of the C1 and the C2 channels. An internal levee is present in the northern side of the C2 channel. A partial depositional setting, consisting of channelwide HARs characterizes the C1 slope channel. c) N–S trending Air Gun seismic proﬁle of the lower segments of the C1 slope channels showing the internal architecture of its internal levee and of the inﬁll of the two slope channel, characterized by High-amplitude Reﬂections (HARs). d) N–S trending Air Gun seismic crossing the distal lower segments of the C1 and C2 channels that are partially inﬁlled by HARs corresponding with conﬁned intrachannel lobes.
5.2. Sedimentary features along the bend sectors Besides the general sedimentary setting of the C1 and C2 slope channels, more complex processes occur in the bend sectors, where a variety of features are recognized. Knickpoint A forms upslope from
the ﬁrst bend of the C1 slope channel, where the channel ﬂank is at 90° to the main ﬂow direction (Figs. 5, 8). We suggest that Knickpoint A formed in response to the reﬂection of sedimentary ﬂows against this ﬂank. As many outcrops example and theoretical and experimental analysis show, the interaction of sedimentary ﬂows against an
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C1 slope channel Inner thalweg
Knickpoints -720 m
Knickpoint A -800 m
8° -920 m
7° 6° 5° 4° 3° 2° 1° 0
C2 slope channel
C2a reach Inner thalweg
Lower reach -840 m -880 m
9° 8° 7° 6° 5° 4°
3° 2° 1° 0
Fig. 7. Longitudinal slope proﬁle (meters below sea level, left) and slope gradient (degrees, right) of the C1 (a) and C2 (b) slope channels with the location of inner thalweg sectors, and of the main knickpoints. The limit between the upper and the lower reaches of the two slope channels is also indicated.
obstacle can generate ﬂow disturbances, that can lead to an upstream migrating bore, able to produce a current ‘reﬂection’ (Kneller, 1995; Kneller and Buckee, 2000). We thus argue that the megascour associated with Knickpoint A can be the result of ﬂow reﬂection against the channel ﬂank. The reﬂected ﬂow experiences a hydraulic jump at the base of the channel ﬂank, and the related increase in turbulence can enhance the excavation of the channel ﬂoor (see Edwards, et al., 1994; Kneller and Buckee, 2000; Satur et al., 2004). A similar morphology, that could be the result of the interaction between channel topography and sediment
gravity ﬂows, is observed between the ﬁrst and the second turns of the C2 channel (Fig. 5). In this case too, corresponding to the turn, a marked knickpoint forms, associated with a narrow, spoon-shaped depression (Figs. 5, 7, 8). The unusual morphology of this sector could be the result of a hydraulic jump and the transition from supercritical to subcritical ﬂow in correspondence with the obstacle effect in the bend. This ﬂow transition results in a loss of ﬂow capacity and waning sedimentary ﬂows downstream from the knickpoint, and could be responsible for enhanced deposition (Hiscott, 1994; Kneller, 1995;
G. Dalla Valle, F. Gamberi / Marine Geology 286 (2011) 95–105
a 810 m
b Channel flank Megascour
A 950 m
970 m 500 m
Fig. 8. 3D shaded reliefs of the bend sectors of the C1 (a) and C2 (b) slope channels, with details of their main geomorphic elements. Below each 3D map, a longitudinal proﬁle of the turns is shown (see the dotted line on the 3D map for the location of the proﬁles).
Kneller and McCaffrey, 1999; Heiniö and Davies, 2007). This behavior can explain the ﬂat morphology of the spoon-shaped depression, that could be the site of the deposition of coarse-grained materials (Garcia and Parker, 1989; Walthman, 2004; Heiniö and Davies, 2007). Channel conﬁned knickpoints can have various origins, but mainly originate where changes in ﬂow properties occur in response to variations in slope gradient, often brought about by allogenic factors such as tectonics (or halokinesis), or in the degree of topographic conﬁnements (Mitchell, 2006; Heiniö and Davies, 2007). In the C1 and C2 cases, the knickpoints formation is mainly driven by the interaction between the sediment-gravity ﬂows and the planform of the channel. The most striking elements of the bend sectors of the C1 and C2 slope channels are erosional features such as knickpoints, megascours and erosional terraces (Fig. 5). Therefore, even if the channels are sinuous, they do not develop true ﬂuvial-style, point bars. There are very few examples of unambiguously truly point bars in the submarine channel literature, and most of them come from very coarse-grained, conglomeratic systems (Dykstra and Kneller, 2009). In other cases, inner bend deposits are small features which are probably lateral accretion associated with smaller intra-channel elements (Deptuck et al., 2007; Kane and Hodgson, 2011). Therefore, it is plausible that slope channel systems characterized by high bedload transport (i.e. gravel-bearing systems) may be more prone than sand or mud dominated systems to develop ﬂuvial-style point bars. Accordingly, the lack of depositional point bars in the bend sectors of the C1 and C2 slope channels, could be caused by the grain size of the turbidity currents ﬂowing within the two channels. Notwithstanding the lack of cores in the study area, taking into account the hinterland geology of northern Sardinia, characterized mainly by Varisican granitoids (Franceschelli et al., 2005), and the grain size of the coastal areas, a sandy character for the ﬂows within the C1 and C2 slope channels is to be expected. In addition, a similar grain size is in agreement with that of the nearby Golo turbidite systems, where ground truthing samples are available, (Deptuck et al., 2008).
5.3. Sea level variations, sediment sources and channel evolution The present-day setting of the C1 and C2, interpreted to reﬂect a phase of reduction of sediment gravity ﬂow energy, matches the recent evolution of the Caprera deep-sea fan as reconstructed by Dalla Valle and Gamberi (2010). The shelf facing the OB slope is around 20 km wide and thus, it can be reasonably assumed that, at present, the heads of the slope channels are distant from the direct sediment input from the short rivers that outﬂow in inlets and small gulfs of the Sardinian coastline (see Ulzega, 1987; Gamberi and Dalla Valle, 2009; Piper and Normark, 2009; Romans et al., 2009) (Fig. 1). However, during low stand periods, when the coastline nearly coincides with the shelf-break, the slope channels were likely to be directly connected with eventual shelf-margin deltas. We thus argue that during sea level low-stand the Caprera slope channels were sectors of prevalent erosion and sediment bypass. Concomitantly, as shown by Dalla Valle and Gamberi (2010), the distal part of the Caprera deep-sea fan was the main depositional site. Although at present, river mouths are far from the channel heads, the erosional channel ﬂoor of C2a, and the inner thalweg within the upper parts of the other slope channels, show that sediment gravity ﬂows are still focussed within the CSCS. A mechanism for sediment gravity ﬂows initiation and downslope sediment delivery must therefore still be active during the presentday high stand. In general, during high-stand periods much of the coarse-grained sediment carried by the rivers is stored on the shelf, and episodic stormrelated turbidity currents, wind-driven and cascading currents, are the only processes capable of sediment transfer to the deep-water (Thornton, 1984; Piper and Normark, 2009). In particular, the interaction between alongshore currents and currents perpendicular to the shelf has been recognized as a possible agent for shelf to slope sediment transfer during high-stand conditions (Paull et al., 2005; Covault et al., 2007; Boyd et al., 2008; Normark et al., 2009a). Likewise, in the CSCS, the main sedimentary input could come from the interaction between the southward directed
G. Dalla Valle, F. Gamberi / Marine Geology 286 (2011) 95–105
littoral cell ﬂowing along the Corsica Island (Astraldi and Gasperini, 1992) and the wind-driven eastward outﬂowing currents from the Bonifacio Strait (Artale et al., 1994) (Fig. 1). 5.4. Channel demise and passive healing A more complex history must characterize the C0 slope channel, that has its upper segment completely healed by the deposition of about 80 ms of sediment pointing to a longer period of inactivity with respect to the C1 and C2. The C0 is the farthest of the CSCS channels from the Bonifacio strait (Fig. 1). What is more, the small V-shaped channel-forms, which develop adjacent to the C0 slope channel, are, at present, buried and healed (Figs. 2, 4a). In general, it is the position of the head of a slope channel that controls the amount of sediments entering the fairway (Covault et al., 2007; Jobe et al., 2011). Thus, it is possible that during high stand periods the effects of the interaction between the Corsica long shore current and the Bonifacio strait outﬂow current have less probability of reaching the head of the C0 slope channels and of the V-shaped channel-forms. We thus speculate that the initial healing of the proximal, upper part of the C0 channel occurred during a past high stand, and led to the detachment of the C0 head from direct river input, which resulted in its early abandonment with respect to C1 and C2. The abandonment of the C0 and the V-shaped channel-forms has taken place through several cycles of eustatic variations, with sedimentary ﬂows that, in this sector of the continental slope are not able to erode or to rejuvenate the sedimentary fairways, but have a predominant depositional character. We argue that, also during low-stand periods diluted, muddy, unconﬁned sedimentary ﬂows predominate in this slope sector. Therefore, they have caused the a progressive healing of the C0 and the burial of the V-shaped channel-forms (see Jobe et al., 2011), ﬁnally resulting in the deposition, of a lobe restricted to the upper slope sector (Fig. 4a). A similar abandonment of a submarine fairway by the progressive healing of its head and uppermost reaches due to high-stand related sedimentation has been documented by Walsh et al. (2007). Jobe et al. (2011)have also reported the passive inﬁlling of canyons by the widespread deposition of mud-prone draping units in the Paleogene of Equatorial Guinea. A comparable passive healing style, seaward from the shelf, has also been documented in the Miocene canyons of the New Jersey margin by Mountain et al. (1996), where the upper segments of the fairways were healed by hemipelagic mud pinching-out downcanyon. 6. Conclusions The sedimentary evolution of the CSCS demonstrates the important role that oceanographic conditions on the shelf play in transferring sediment to the heads of slope channels. It shows that in a margin lacking in major rivers, and with a wide shelf, during sea level high stand the interaction between longshore and cross-shelf currents determine downslope sediment delivery and the degree of slope channel activity. The CSCS also illustrates the response of slope channels to the decrease of sediment gravity ﬂow volume associated with allogenic factors control such as sea level variations. During sea level low-stand conditions, large sediment gravity ﬂows, presumably river-connected, are capable of eroding wide slope sediment fairways that can reach the base of the slope to feed a basin plain fan. During the present-day highstand of sea level, the slope channels are not completely abandoned, but are the site of small sediment gravity ﬂows. The upper reaches of the slope channels are in fact characterized by erosional and bypass processes with the formation of inner thalwegs and erosional terraces and by intrachannel depositional processes that result in internal levee construction. On the contrary, their lower reaches are the site of prevailing backﬁlling processes with the aggradation of channel-wide HARs which are conﬁned within the master channel ﬂanks.
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